ES 10
Lecture 11 - Global Seismology
Greg Anderson

Reading: Blue Planet Chapter 3

``I'm afraid I can't comment on the name Rain God at this present time,
and we are calling him an example of a Spontaneous Para-Causal
Meteorological Phenomenon.''

``Can you tell us what that means?''

``I'm not altogether sure. Let's be straight here. If we find something
we can't understand we like to call it something you can't understand, or
indeed pronounce...

``...And if it turns out that you're right, you'll still be wrong, because
we will simply call him a ... er, `Supernormal' -- not paranormal or
supernatural because you think you know what those mean now, no, a
`Supernormal Incremental Precipitation Inducer.' We'll probably want to
shove a `Quasi' in there somewhere to protect ourselves. Rain God! Huh,
never heard such nonsense in my life. Admittedly, you wouldn't catch me
going on holiday with him. Thanks, that'll be all for now, other than to
say `Hi!' to Wonko if he's watching.''

Central Ideas
Basic Terms and Questions

Before you can learn anything about global, regional, or local seismology,
there are a few basic ideas you need to have seen. We begin with some
basic terms:

Seismology

The scientific study of earthquakes.

Seismologist

A scientist who studies earthquakes.

Seismometer

An instrument which records earthquakes.

Seismograph

A device which takes the signals recorded by a seismometer and makes
a recording (on paper, on film, in a computer...) of those signals.

Seismogram

The recording made by a seismograph of the signals from a
seismometer.

Fault

A crack in the ground across which there is relative movement
between two adjoining blocks of rock. It is the movement
which makes a fault a fault --- otherwise, it's just a crack or
a joint. No, not that kind of crack or joint.
DeKalb College in Georgia has a
good simple page
on different types of faults and Lisa Tauxe will talk about faults
more later in the class.

What is an earthquake?

What exactly is an earthquake? An earthquake is a natural
event inside the Earth which releases built-up energy. Earthquakes
release energy and create seismic waves, which cause the shaking we
feel. It's important to be clear: an earthquake is not the
shaking we feel. An earthquake is the event which makes
the seismic waves (we'll talk about them in a bit) which cause the
shaking we feel.

How and why do earthquakes happen?

In 1906, there was a large earthquake along the San Andreas Fault
in northern California. In the course of field investigations,
which among other surprises showed that some spots along the San
Andreas had moved by as much as 7 meters, H.F. Reid developed the
theory which is most used today to explain earthquakes: the
elastic rebound theory.

The elastic rebound theory basically works like this. The rocks
along a fault are held together by friction and cannot slip relative
to each other. Over time, background geologic forces (which we now
know to be mostly plate motions) apply a stress to the rocks on
either side of the fault. The blocks of rock ``want'' to slide
past one another, but cannot because of the friction holding them
together. Instead, they deform around the fault, building up
strain energy. Eventually, there comes a point when the strain
energy is greater than the strength of the friction holding the
rocks together, and they slide past one another abruptly.
Voilà -- an earthquake.

There are other kinds of earthquakes which are not directly caused
by faults sliding due to plate tectonic stresses. Earthquakes are
sometimes caused by volcanoes, by filling reservoirs, by pumping
fluids in or out of the ground, and by stresses imposed by big
underground nuclear explosions. These sorts of earthquakes are
certainly the minority, however.

Central Ideas
Seismic Waves

When an earthquake occurs, it makes seismic waves, which cause the shaking we
feel. Seismic waves are essentially just the jiggling of
the ground in response to the force put on the ground by the earthquake,
similar to the way the jello in a bowl responds to a tap to the side of the
bowl. There are three major kinds of seismic waves: P, S, and surface
waves. P and S waves together are sometimes called body waves
because they can travel through the body of the earth, and are not trapped
near the surface.

A P wave is a sound wave traveling through rock. In a P wave,
the rock particles are alternately squished together and pulled apart
(called compressions and dilatations), so P waves are also
called compressional waves. These waves can travel through solids, liquids,
and gases. P waves can travel through the liquid outer core.

An S wave is a different beast. In an S wave, the rock
particles slide past
one another, undergoing shear -- so an S wave is also called a
shear wave. You can make shear waves by, for example, tying a rope to a
tree and shaking the free end of the rope up and down or side-to-side.
The waves themselves will travel forward, toward the tree. But the rope
particles will stay in one place, sliding back and forth past each other.
Shear waves cannot travel in liquids or gases -- so, for example, S waves
don't travel through the ocean or through the outer core.

Surface waves are called surface waves because they are
trapped near the
Earth's surface, rather than traveling through the ``body'' of the earth
like P and S waves. There are two major kinds of surface waves: Love waves,
which are shear waves trapped near the surface, and Rayleigh waves, which have
rock particle motions that are very similar to the motions of water particles
in ocean waves.

Keep in mind that P- and S-waves travel at different speeds, but that both
P- and S-wave speeds depend on density and ``stiffness'' of the rock they
travel through. Here are the equations which govern wave speed:

What you need to know about this is that

P- and S- waves travel at different speeds,

Wave speed goes up as material gets stiffer

Wave speed goes down as material gets denser

Also, note that seismic waves can be reflected or refracted at
places where the rock properties change, and that these reflections and
refractions can be used to study the earth's structure. Figures 3.6 and
3.7 in The Blue Planet show reflection and refraction of
seismic waves.

Central Ideas
How do seismologists record earthquakes?

Seismology is the science of studying earthquakes.
Seismologists are scientists who study earthquakes.
We record ground shaking with an instrument called a seismometer, and
the instrument makes a recording on a device called a seismograph --
sometimes on paper with ink, but mostly these days with digital computers. The
recording itself is called a seismogram. This is, of course, all
review from earlier, but it helps introduce the next bit...

Most classical seismometers have either a heavy mass on a suspension system,
like a spring, or a mass at the end of an arm which swings like a fence gate.
Seismometers work by sensing the relative
motion of the heavy mass and the frame of the seismometer itself. The mass
``wants'' to stay in one place due to inertia, while the frame of the
instrument has to move with the ground, since the frame is firmly attached to
the ground. This relative motion is sensed by the instrument, and is what is
recorded by the seismograph to make a seismogram.

Most modern instruments are actually completely computerized, and work by
sensing how hard they have to work to make the mass move with the
rest of
the instrument. This record of the force necessary to make the mass move is
stored digitally in a computer connected to the seismometer, and sent via phone
lines or the Internet to a processing center, where seismologists use computers
to look at the records and play with the earthquakes. These days, most
seismic data processing never actually involves paper records -- though I have
one of my favorite paper records in a frame on my wall.

Below are a few example seismograms which I made using the UC Berkeley tool.
Most are seismograms from the CMB station, which is in the Sierra foothills
at Columbia College. The Northridge record is from the BKS site which is
located just east of the UC Berkeley campus.

This is a recording of a magnitude 3.6 earthquake 13 km SE of Mammoth Lakes
which happened at 1626 UT (8:26 am PST) on 24 February 1997. Note the clear P
and S wave arrivals at about 16:26:18 and 16:26:35 respectively.

This is a recording of the 17 January 1994 Northridge earthquake (M6.7) near
Los Angeles at 1230 UT (5:30 AM PST). This earthquake is about 525 km away
from BKS. Note there is an M 5.9 aftershock which is lost in the arrivals
from the main event. Other aftershocks are visible at 1507 (M 4.2), 2046
(M 4.9), and 2333 UT (M 5.6).

This is a recording of a magnitude 6.5 earthquake in the Andreanof Islands on
08 June 1996 at 2319 UT (4:19 pm PDT). This earthquake is about 4700 km away
from CMB and is a shallow earthquake. The waves with the highest
amplitude (the ``biggest wiggles'') are the surface waves, which is typical
in large, shallow, distant earthquakes.

This is a recording of a magnitude 6.9 earthquake in the Tonga Islands on 19
October 1996 at 1453 UT (6:53 am PST). This earthquake is about 8900 km from
CMB and is a deep earthquake, with a depth of about 590 kilometers. Note that
the surface waves here are fairly weak; typically, deep earthquakes don't
generate strong surface waves.

One question lots of people ask me is: ``How do you locate an earthquake?''
It turns out that, while the procedure is not entirely straightforward, it is
not all that difficult to locate an earthquake. I'll tell you how we locate
local earthquakes; distant earthquakes are located using similar methods, but
they are a bit more complicated than we need to worry about.

You will recall from the discussion earlier that there
are three major kinds of seismic waves: P, S, and surface waves. P waves
travel faster through the Earth than do S waves, so P waves arrive before S
waves do. If you have a seismogram, and you know how to measure time
accurately on it, you can pick the arrival times of the P wave and
the S wave. Next, you figure out how far apart these waves arrive, called
the S-P time. You can then go to a table of distance as a function
of S-P time and work out how far away the earthquake was from your station.
If you have three or more stations, you can draw circles on a map, and where
the circles meet is the location of your earthquake. Essentially, you are
triangulating the earthquake's location.

An example is probably useful. Below, you will see a seismogram
from a seismometer in northern California called BKS. You should see an
earthquake recording (the squiggly line), and specifically note that there are
two points in the record
at which the size of the signal (what seismologists call the amplitude)
jumps fairly strongly. The first jump is the arrival of the P wave, and
the second major jump (although in this case, it's actually a major drop) is
the arrival of the S wave.

Based on the time scale shown on the plot, I can figure out the S-P time for
this earthquake recorded at this station. In this case, the P wave arrives
at 11.2 seconds after 04:33, and the S wave arrives at 15.8 seconds after
04:33. My S-P time is 4.6 seconds. Now, given a table which tells me how far
away the earthquake is if I can tell the time difference, I find that the
earthquake is about 36 km from BKS. (If you ever have a homework problem on
this subject, you certainly can expect me to give you such a table, even though
I don't show one here...)

Below are seismograms from three more sites in northern California, BRIB,
JRSC, and MHC.

Following the
same procedures to get S-P time and distance for these stations that I did for
BKS, I can make the following table:

Table 1: Location Information Example

Now that I have my distances, I can get out my map of northern California,
locate the stations on it, and draw circles of the appropriate radius (given
the correct map scale, of course).
My estimate of the location of the earthquake is the point where the circles
intersect. I have marked the location that UC Berkeley finally settled on
(using much more data than I have) with an inverted triangle -- you can see
that the two locations are not much different.

Of course, these days nobody really locates earthquakes as I have just shown.
Earthquakes are now located using computers and models of the structure of
the Earth's crust in a given region. With digital recording, a skilled human
operator, and good models of the crustal structure, locations can be found
which are accurate to within about 250-500 meters.

Note that the US Geological Survey's National Earthquake Information
Center has a
very nice web page on different ways to measure earthquakes.
I certainly recommend reading it in addition to what I have to say below!

Most people living in California have heard about the ``Richter Scale'' and
have at least a vague idea that it is used to measure the sizes of earthquakes.
Most people that I know, though, have some misconceptions about the Richter
Scale -- for example, someone once asked me if he could see where I kept
my Richter Scale! Also, there are some important differences between
magnitude, energy, and intensity that need to be discussed.

C.F. Richter at Caltech invented the idea of earthquake magnitudes in 1935 as
a way to compare earthquakes. He was into astronomy and knew that astronomers
used magnitude scales to compare the brightnesses of stars, so he adapted the
idea for seismology. Richter based his scale on the way ground motion was
recorded by a specific type of seismometer that was very common back in the
1930's, called a Wood-Anderson seismometer. It is very important for you to
realize here that the Richter Scale is completely arbitrary; it was
made up by Richter.

Basically, if you know how far away an earthquake is from your station, and you
have a record from the earthquake, you can calculate its Richter magnitude.
You do that by measuring the maximum amplitude of the shaking recorded
by the W-A instrument, taking the logarithm of that amplitude, and adding in a
number to take into account distance from the station. The key thing here is
that the magnitude scale is logarithmic; for every one full point
change in magnitude, the amount of shaking recorded by a seismometer will go
up by a factor of 10.

Richter magnitude is useful, but limited; it is only defined for local
earthquakes. Other magnitude scales have been developed to handle earthquakes
that are distant from the seismometer making a given magnitude estimate, and
these have been made to give magnitudes which are basically similar to Richter
magnitudes. However, all of these scales have the same two fatal flaws.
First, all of them become inaccurate at large magnitudes, and in fact, above
about magnitude 8 or so, the magnitudes just don't get bigger (even though
the earthquakes do). Second, these magnitude scales are all empirical; they
don't actually tell you anything about the physics of the earthquake itself.

Recently, a new magnitude scale has been developed which has neither of these
flaws: the moment magnitude scale. Unlike all the other kinds of
magnitude, moment magnitude tells you something about the physical size of
the earthquake. Moment magnitude really is physically meaningful. Also,
moment magnitude is never overwhelmed by large earthquakes, and so it is
possible to get meaningful estimates of magnitude even for humungous quakes.
As an example of this, recently people have gone back and recomputed magnitudes
for earthquakes such as the 1964 Alaska earthquake. Previously, the magnitude
had been given as 8.6 - pretty damn big, but nothing compared to the 9.2 which
is now accepted (remember that magnitude scales are logarithmic).
Moment magnitudes are now the accepted magnitude among seismologists, and are
usually the numbers given to the press.

Another way of looking at the size of earthquakes is to figure out how much
energy they release. Some rules of thumb have been found to compare
magnitude to energy, and it has been found that for each one point magnitude
increase (say from a 4 to a 5), 32 times as much energy is released. If one
jumps from a 5 to a 8, the energy goes up by 32 x 32 x 32,
which is almost a factor of 33,000 -- but don't worry. While the total energy
goes up that much, it does so not because the ground shakes 33,000 times
harder, but instead because large earthquakes release energy for much longer
and over a much wider area than do smaller earthquakes.

Finally, there is another way of looking at the strength of earthquakes, which
depends not on records of earthquakes but on how the earthquake was perceived
by people and how much damage is done. This is called intensity, and
is described using the so-called
Modified Mercalli Scale. (As a sidelight, Roger Musson from the
Global
Seismology and Geomagnetism Group of the British Geological Survey
has assembled an
interesting history of
the "Mercalli Scale", mainly saying why it's not really the Modified
Mercalli Scale).
While there will be very little
variation in magnitude estimates for a given earthquake, intensity measurements
can (and do) vary widely. Intensities vary based on the distance from the
earthquake, what the person making an intensity report was doing at the time
of the earthquake (intensity would be lower from someone who was air-guitaring
to Van Halen than from someone sitting quietly playing a viola, for example),
by what kind of building they were in, what kind of soil they were on, etc.
Intensities are inherently subjective, but can be of use to engineers who try
to build earthquake-resistant buildings, for example.

OK, so how big do they get?

How big can an earthquake get? Unfortunately, I can't give you a simple answer
here.
The maximum magnitude earthquake that a given fault can generate is determined
by a number of factors. A long fault can generate a larger earthquake than
a short fault, all other things being equal. A fault which tends to break and
have the rocks jump farther than another fault will tend to generate larger
earthquakes than that other fault. And the strength of the rocks is another
factor: stronger rocks will tend to hold out longer, and generally break in
a larger earthquake, than will weaker rocks.

I'm not sure how big the theoretical maximum size earthquake is period, but I
can tell you that
the largest earthquake ever recorded was in Chile, on 22 May 1960. It had
a moment magnitude of 9.6, broke an area of fault 850 kilometers long and
more than 120 kilometers wide, and generated a lot of damage and a humungous
(taller than 30 feet in some places along the Chilean coast) tsunami. This
earthquake released about as much energy as would be released by blowing up
one billion tons of TNT!

At the other end of things, there is no limit to how small earthquakes can
get. The instrumental limit is that, in the quietest locations with the most
sensitive seismometers and earthquakes extremely near to the seismometer, it is
possible to record earthquakes as small as magnitude -2. An earthquake that
small would rupture a circular fault roughly 7 centimeters across and move it
about 1 centimeter -- a tiny, tiny, tiny earthquake.

People generally stop feeling earthquakes when they drop below about magnitude
3 or so, although I know of a case where a magnitude 2.3 earthquake was felt
by someone sitting very quietly in a house which was right on top of the
epicenter.

Just as an aside, there was a great earthquake (magnitude 8.2) in Bolivia
in 1994 at a depth of about 630 kilometers which was actually felt in
North America. The earthquake was felt in high-rise buildings as
far away as Renton, Washington -- which is almost 8700 kilometers from
the epicenter! This is the greatest distance over which any earthquake is
known to have been felt anywhere in the world.

Central Ideas
How often do they happen?

Earthquakes happen constantly around the world. In an average year, there
might be 20-25 magnitude 7 earthquakes globally -- about one every 2 to
3 weeks. On the other end of the magnitude scale,
there are literally hundreds of thousands of tiny earthquakes worldwide in a
given year.

Locations of most earthquakes

More than 95% of the world's earthquakes occur in discrete belts throughout
the world. The existence of these belts is one piece of evidence in support
of plate tectonics. In fact, we now know these belts are plate boundaries.
Earthquakes occur along all types of plate boundaries: subduction zones,
transform faults, and spreading centers.

However, there are earthquakes which occur within plates. For example, the
New Madrid area of Missouri (1811-1812), Charleston, South Carolina (1889),
Boston (1755), and Hawaii (1975) are all places which have large earthquakes.
In fact, in 1811-12 there were
four very large earthquakes in the New Madrid area, which are believed to
have been in the low magnitude 8 range (there are no actual recordings from
which to figure out magnitudes...). These earthquakes actually rang church
bells in Boston -- over 1700 kilometers away! -- and caused damage as far
away as Washington D.C, more than 1100 kilometers away. There are also
stories that say these earthquakes made the Mississippi River flow
backwards for a short time!

Earthquakes which occur within plates (intraplate earthquakes) are
among the remaining mysteries for plate tectonics, because plate tectonics
cannot strictly explain their occurrence. In some places, such as Hawaii, the
earthquakes are related to volcanism, but for the most part, intraplate
earthquakes are as yet not fully understood.

Here's a map of earthquakes globally over the past five years, generated
using the
Seismic Monitor
tool from the Incorporated Research Institutions for Seismology (IRIS).
As you can see, most earthquakes are at plate boundaries, but there are
some intraplate earthquakes. This map is up to date as of 2 February 1997
4:04:13 PM PST.

To generate a new map for yourself, use this
web page and click on the Seismic Monitor button. The
new map will be at most 30 minutes old.

Why can moderate depth or deep earthquakes exist?

In the late 1920s, scientists first were able to prove that earthquakes
happened in regions where were parallel to deep ocean trenches, inclined
about 40-60° from the horizontal, and extended to great depths within
the earth (several hundred kilometers). The two seismologists most responsible
for recognizing these belts were Kiyoo Wadati in Japan and Hugo Benioff in
the USA, and over time, they have come to be called
Wadati-Benioff zones or Benioff zones.

These regions were known certainly by the early 1930's, but were not explained
by then current ideas. It was only after the development of plate tectonics,
which stated that plates were subducting into the mantle at deep trenches,
that these zones could be explained. Keep in mind that these earthquakes are
happening at depths far beyond the maximum depth at which the rocks are cool
enough to support earthquake fractures. During the development of plate
tectonics, deep and moderate-depth earthquakes were found to be located inside
subducting slabs, which are much cooler than the rocks around them and can
support rupture during earthquakes.

It is worth noting, however, that the cause of deep earthquakes is
still being debated even today.

Earthquakes as a Tool

Earthquakes generate seismic waves, which can travel great distances
through the earth. From these seismic waves, scientists can infer the
structure of the earth on
a global scale. We discussed earth structure in previous lectures, so
I won't dwell on it here. I will simply give you a brief listing
of some of what I think are the most interesting bits in global seismology.

You know the earth is divided into (basically) the crust, mantle, outer and
inner core, like the following figure:

We infer this structure from many thousands of measurements of the time it
takes seismic waves to travel around the world from various earthquakes.
These data can tell us the seismic wave velocity and (remember the formulae
for wave speed) the density at various points in the earth, and other data
are combined with that to make the model we have of the inside of the Earth.
Here's another model, which shows velocity and density as well. (Thanks
to Peter Shearer for this figure.)

OK, so the basic radial earth structure is known. Now the most active parts
of research lie in looking for variations in the properties of the Earth in
all three dimensions, not just in distance out from the center. Among the
areas which draw lots of attention are

The upper mantle, which exhibits various changes in mineral structure
which result in changes in the seismic velocity at depths where those
mineral changes happen. These are known as upper mantle
discontinuities, and lots of people around the world study them,
including some folks at SIO. They may play an important role in
controlling subduction and convection on Earth.

The lowermost mantle (called D'' by seismologists and others), which
is the largest change in chemical composition and seismic velocity and
density inside Earth. At D'', we switch from a silicate rock-based
mantle to a liquid iron-nickel-etc based outer core (movement of which
generates the geomagnetic field). Lots of weird stuff is going on
at the core-mantle boundary and D'', most of which is not understood.

The inner core, which is largely solidish iron/nickel/other gunk and
forms by cooling of the outer core. The outer part of the inner core
appears to be sort of mushy, and the rest of it is largely solid.
The whole thing may be
spinning at a faster rate than the rest of the planet. This is
really an interesting result, and we aren't sure if it's right, but
it's one of the hot topics lately...